"Chemical fossil" was defined as an organic com- pound whose carbon skeleton suggested an unam- biguous link with a known natural product (Eglinton and Calvin, 1967). Al though the term "chemical fossil" has been replaced by the term "biological marker" or "b iomarker" in recent years (Philp, 1985; Johns, 1986; Engel and Macko, 1993), the usefulness of this idea appears to have increased year by year, especially during the last decade, as a tool for reconstituting paleoenvironments. Studies applying biomarkers to the reconstruction of paleoenvironments have often been called "molecu- lar paleontology". Compared with "ord inary" paleontology, however, molecular paleontology has several advantages.
First, molecular paleontology involves infor- mation of organisms which do not have hard parts, such as carbonates or silicates. Because organisms, such as bacteria or blue-green algae which do not precipitate hard parts have played a major role in the chemical evolution of biosphere during the last
*Present address: Institute of Low Temperature, Hokkaido University, NI9W8 Kita-ku Sapporo 060, Japan.
4 billion years, direct information on these organ- isms should contribute to progress in understanding developments of geochemical processes and cycles. Secondly, the relative abundances of several com- pounds are known as a function of environmental forcing, such as temperature or pH when they are biosynthesized (Harwood and Russell, 1984). Therefore, quantitative reconstruction of these par- ameters can be attained once the calibration is determined. Thirdly, the evolutionary rate of bio- chemical processes, such as carbon assimilation is much slower than that of the morphology in organ- isms, which allows us to directly extrapolate knowl- edge on modern biochemistry to organisms that are already extinct.
The paleoenvironments of the late Quaternary, characterized by glacial-interglacial climatic vari- abilities, have been studied intensively with empha- sis on carbon cycles during the last decade (e.g. Zahn and Pedersen, 1994) in order to elucidate the secular variations of carbon dioxide concentration in the atmosphere during that time (Barnola et al., 1987; Neftel et al., 1988). The biomarker appli- cations to the late Quaternary paleoclimatology were pioneered by Brassell et al. (1986). They pro-
173
174 N. Ohkouch i et al.
posed the unsaturation degree of C37 alkenones as a useful stratigraphic tool, which has been fol- lowed by many studies (e.g. ten Haven et al., 1987; Jasper and Gagosian, 1989; Eglinton et al., 1992; Kennedy and Brassell, 1992; Rostek et al., 1993). Prahl et al. (1989) first analyzed the carbon- ate-rich sediments collected from the central tropi- cal Pacific for free- and bound-form fatty acids during the last 26 000 y. They estimated that about 20% of total organic carbon preserved in the sedi- ments originated in terrestrial organisms and was introduced to the core site by long-range atmos- pheric transport. With subsequent results and in- terpretation on biomarker abundances in this core, Prahl (1992) suggested that the biological pro- duction in the central tropical Pacific could have been controlled by atmospheric iron flux during this time-span.
To use organic compounds in deep-sea sediments as quantitative tools for reconstituting past environ- mental factors, calibrations between accumulating organic compounds and oceanographic and/or cli- matological factors should be determined rigidly by the observations of today's ocean environments. We have reported lipid class compounds in the pelagic sediments from the Pacific Ocean (Ohkouchi et al., 1997b), and the box core sediments from the wes- tern tropical Pacific (Ohkouchi et al., 1994; Ohkouchi et al., 1997a). In this paper we combine the lipid data and discuss climatological changes from the last glacial maximum to the present with consideration of the principles for deducing three environmental factors: terrestrial input to the ocean environments; marine biological productivity; and sea surface temperature.
EXPERIMENTAL
Samples
The sampling locations of the sediments used are shown in Fig. 1 (Ohkouchi, 1995; Ohkouchi et al., 1997a, 1997b). Twenty-three deep-sea surface sedi- ments were recovered from the central Pacific with a box corer (0-2 or 0-4 cm) located on the longi- tude of 175°E and spread from 48°N to 15°S. A 30 cm-long box core (KH92-1-5cBX) was recovered from the western tropical Pacific (3°32'N, 141°52'E, water depth of 2282 m). The core was cut at 2-cm intervals into 15 time-series sections. The chronol- ogy of this core was determined based on accelera- tor mass spectrometer ~4C dating of the planktonic foraminifera Pleniat ina spp. and graphic correlation to the standard oxygen isotopic curve (Ikehara, 1994; Ohkouchi, 1995). The sediment samples were stored in a refrigerator or a freezer until analysis.
Analys i s
To isolate lipid class compounds, wet sediments were extracted with methanol, CH2C12/methanol
(7:1) and CH2C12:methanol (10:1) using both hom- ogenizer and ultrasonicator. Extracts were washed with 0.15 M HC1 to remove salts contained in the interstitial water of the sediments and saponified with 0.5 M KOH-methanol under a reflux. Neutral lipids were separated from the extracts by extrac- tion with CH2C12:hexane (10:1). They were further separated into four sub-fractions on a Pasteur pip- ette column packed with silica gel (BIO-SIL A, 200 400 mesh) which was deactivated with 1% water. Aliphatic hydrocarbons (N-l), polynuclear aromatic hydrocarbons (N-2), ketones and alde- hydes (N-3) and fatty alcohols and sterols (N-4) were eluted with n-hexane, n-hexane:CHzC12 (2:1), CH2C12 and CH2Clz:methanol (95:5), respectively. The N-4 fraction was treated with bis-trimethyl- silyl-trifluoroacetamide (BSTFA) prior to gas chro- matographic analysis. Acidic lipids were separated from the extracts with CH2C12 after the remaining solution was acidified to < pH 2. They were deriva- tized to methyl esters with 14% BF3/methanol at 100°C for 30 min. The acidic lipids separated into three sub-fractions on the silica gel column: mono- carboxylic acid methyl esters (A-l); dicarboxylic acid and ketoacid methyl esters (A-2); and hydroxy acid methyl esters (A-3) by eluting with n- hexane:CHzCl2 (1:2), CH2C12:ethylacetate (98:2) and CH2Clz:methanol (95:5), respectively.
Gas chromatographic analyses were performed with Carlo Erba 6000 gas chromatograph (GC) installed with an on-column injector, HP-5 fused silica capillary column (25 m x 0 . 3 2 m m i.d.; 0.52pm film thickness), and a flame ionization detector (FID). N-3 fraction was analyzed with Carlo Erba 5160 GC, which was installed with DB- 5 (30 m × 0.32 mm i.d.; 0.25/tm film thickness). Hydrogen was used as a carrier gas and the column oven temperature was programmed from 70°C (1 rain) to 320°C (30rain) at 6°C/rain. The GC peaks were processed using a Shimadzu Chromatopac C-R7A integrator.
Identification of each peak was conducted by comparing gas chromatographic retention times and mass spectra with those of authentic standards and/ or those published in the literature (e.g. Budzikiewicz et al., 1967; Philp, 1985). Gas chroma- tography-mass spectrometry (GC MS) analyses were performed with Finnigan MAT ITS-40 installed with HP-5 fused silica capillary column (30 m × 0.25 mm i.d.; 0.25/~m film thickness).
RESULTS AND DISCUSSION
Terrestr ial input to the ocean environments
Recently, the atmosphere has been considered to be a more important pathway for material transport from the continent to the open ocean than was pre- viously thought (e.g. Uematsu et al., 1985; Duce, 1989). Significant amounts of mineral dust originat-
Climate variabilities in the late Quaternary 175
0 °
Fig. 1. The locality map of sediments discussed in this study, e: Surface sediments recovered at a 175°E transect from 48°N to 15°S; I : box core KH92-1-5cBX (3°32'N, 141°52'E, water depth of 2282 m); El:
locations of SEAREX air sampling stations (Duce, 1989).
ing from sandstorms in the deserts of East Asia and the Sahara were detected in the atmosphere over the central Pacific and tropical Atlantic, respectively (e.g. Duce et al., 1980; Graccum and Prospero, 1980; Prospero, 1981; Betzer et al., 1988; Merrill, 1989). Numerous organic compounds produced by land plants were also observed in the remote marine atmosphere (e.g. Simoneit, 1978; Gagosian and Peltzer, 1986; Kawamura, 1995), suggesting that they have been deposited on the sea floor as sedi- mentary organic matter.
Because higher molecular weight (HMW) n- alkanes (C25-C35) are one of the major compound groups synthesized as epicuticular waxes of terres- trial higher plants (Eglinton and Hamilton, 1967; Tulloch, 1976), the HMW n-alkanes preserved in oceanic and lacustrine sediments have been regarded as terrestrial biomarkers (e.g. Simoneit, 1977; Kawamura and Ishiwatari, 1985; Brassell, 1993). Figure 2(a) shows the latitudinal distri- butions of HMW n-alkane concentrations in deep- sea surface sediments recovered from the central
Pacific across a latitudinal transect at 175°E. They exhibit significantly high values in the high to mid latitudes (48-25°N) and relatively low values in the low latitudes (13°N-15°S). Obviously, the latitudi- nal distribution of total HMW n-alkanes preserved in the pelagic sediments is similar to that of eolian dust concentrations observed over the Pacific Ocean (Fig. 2(b); Uematsu et al., 1985). The good corre- lation between them confirms that the HMW n- alkanes in the pelagic sediments are derived from the same source regions, i.e. arid areas of East Asia. Therefore, HMW n-alkanes preserved in the pelagic downcore sediments should record vari- ations in the past atmospheric transport of terres- trial materials over the core site.
We analyzed downcore sediments (KH92-1- 5cBX) recovered from the western tropical Pacific for HMW n-alkanes (Fig. 3). The mass accumu- lation rate (MAR) of HMW n-alkanes exhibits a slight decrease from the LGM to the latest period of deglaciation and a rapid increase at the begin- ning of the Holocene. The downcore variations
C25-C35 n-alkanes Mean mineral dust concentrations (p.g/g CaCO3 free dry sed.) (tzg/m3)
Fig. 2. (a) Latitudinal concentrations of total higher molecular weight (C25 C35) n-alkanes normalized b y C a C O 3 free dry wt sediment. (b) The latitudinal distributions of mineral dust concentrations in the modern atmosphere observed at remote islands in the central Pacific (modified from Uematsu et al., 1985). The locations of these islands are shown in Fig. 1. Bars indicate the range of seasonal variations.
Midway Oahu
Enewetak
Fanning Nauru
Funafu t i American Samoa
should depend on two factors: their flux to the ocean floor; and past depositional diagenesis. To know the former, the latter should be removed properly from the profile. Here, an exponential diagenetic model is constructed to eliminate de- composition effects from the downcore profile (Berner, 1980), assuming that the MAR of total HMW n-alkanes was constant during the Holocene (after deglaciation period: 0-7000 y BP) and that decomposition rate has not varied significantly during the last 20000y (Fig. 4a). This model suggests that the atmospheric input at the LGM is twice that of the Holocene and, in turn, it was half that of the Holocene during the latest period of deglaciation (Fig. 4b), although we cannot identify exactly whether these variabilities were caused by changes in the dust generation rate, the pathway of dust transport or the dust scavenging rate from the atmosphere.
There are few applications of terrestrially-derived biomarkers to the reconstitution of aerosol concen- tration during the late Quaternary. Prahl et al. (1989) analyzed a sediment core recovered from the
2 0 20 40 60 80
MAR of C25 C35 n-alkanes (ng/cm2/kyr)
Fig. 3. Downcore mass accumulation rates of terrestrially- derived compounds, total C25-C35 n-alkanes in KH92-1- 5cBX as a function of 14C age. Shaded area indicates the
deglaciation period.
central tropical Pacific (MANOP, site C) for HMW fatty acids to reconstruct the variations of terrestrial input during the last 26 000 y. They found higher concentrations of HMW fatty acids at the LGM and concluded that terrestrial input through long- range atmospheric transport was enhanced during the LGM. In the eastern tropical Atlantic, Poynter et al. (1989) reported that the fluxes of HMW n- alkanes and HMW n-alkanols increased during cold stages, which led them to the conclusion that the trade wind was intensified during the glacial time. In the Japan Sea, Ishiwatari et al. (1994) showed that the concentration of HMW n-alkanes (C23 C35) increased more than three times at the LGM relative to the Holocene. They concluded that the eolian dust load around Japan was significantly higher during the glacial period compared with the interglacial period.
As described above, biomarker results suggest that the atmosphere of the last glacial period was dustier than the Holocene, at least in the western Pacific. Other windblown dusts contained in the pelagic sediments, such as silicate minerals, also suggest a significant increase in terrestrial input over the tropical Atlantic and Pacific during the last glacial period (e.g. Sarnthein et al., 1981; Janecek and Rea, 1985; Hooghiemstra et al., 1987; Rea, 1994).
M a r i n e biological produc t iv i t y
Determination of biological productivity in the past ocean has long been a challenge in paleoceano- graphy, because biological assimilation of carbon plays an important role in the partition of carbon dioxide between the atmosphere and ocean during the late Quaternary (e.g. Martin, 1990). Source- specific biomolecules preserved in the sediments should provide a clue to determination. For example, long-chain (C37 C39) alkenones are bio- synthesized by the class Prymnesiophycae, which includes Emil iania hux ley i and Gephyrocapsa ocea-
Climate variabilities in the late Quaternary 177
0
.
10.
< 15.
20.
<
0 0
0
0 (a)
' ' I ' ' ' I ' ' ' I '
20 40 60
C25-C35 n-alkanes (ng/kyr/cm2)
0
10
15
20
25 , 25 0 80 0 0.5 1 1.5 2 2.5
Relat ive terrestrial input
Fig. 4. Reconstruction of atmospheric input during the last 20 kyr. (a) A thick line indicates an expo- nential diagenetic model to subtract degradation effect from the downcore profile of C2~-C35 n-alkanes in KH92-1-5cBX ((3). The diagenetic model assumes that the terrestrial input during the Holocene (0 7 kyr) did not vary significantly and that the decomposition rate of these compounds has remained con- stant during the last 20 kyr. (b) The relative flux of terrestrial input to the ocean during the last 20 kyr. It is shown as the ratio of analytical value to the model. Shaded area indicates the deglaciation period.
nica as dominant species (Volkman et al., 1980; Marlowe et al., 1984a, 1990). Dinosterol is believed to be a specific marker of dinofragellate (Boon et aL, 1979).
Lower molecular weight (LMW) n-alkanes and pristane could also be used as marine biomarkers under certain environments, al though they are, in a strict sense, multi-source and less specific (Tissot and Welte, 1984). The distributions of L M W n- alkanes (C~7-C20) and pristane across a latitudinal transect at 175°E exhibit two maxima around 48 30°N and 10°N-10°S, with the latter being rela- tively smaller (Fig. 5). In the latitudinal range between 27 and 15°N the concentrations of these compounds are quite low. Their latitudinal distri- butions appear to be reflected by varieties of nutri- ent concentrations in surface water and biological
production in the present surface ocean, which is high in the high latitudes and around the Equator
and is typically low around the subtropical gyre (Fig. 5; Broecker and Sarmiento, 1980; Berger,
1989). It suggests that these compounds can be used as marine biomarkers to reconstruct the past
biological productivity in the surface ocean. However, long-chain alkenones were not detected at a significant level in most surface sediments from
the central Pacific, al though in the present ocean the prymnesiophyte algae largely contribute to the
biological production in the Pacific Ocean (Okada and Honjo, 1973). This may be associated with car-
bonate dissolution, because most sediments were
collected at depths deeper than calcite compensation depth (CCD).
Pl-istane (ng/g CaC03 free dry sed.)
0 5 I0 15 50°N
i 30 .~
20 20 .~
1 0
20°S . . . . i . . . . ~ . . . . ~ . . . . ~ . . . . i . . . . ~ . . . . i i " ' i ' ' ' ~ ' ' ' ~ ' , , i , , ' t ' , , 2 0 ° S
Fig. 5. Latitudinal distributions of marine derived compounds, C~7-C2o n-alkanes and pristane normal- ized by CaCO3 free dry wt sediments and nitrate concentrations in surface waters obtained by GEOSECS (170°E-176°W; Broecker and Sarmiento, 1980). Shaded areas indicate the regions of high
biological productivity in high latitude in the Northern Hemisphere and around the Equator.
Fig. 6. Downcore mass accumulation rates of (upper) total C17 C20 n-alkanes and pristane and (lower) total organic carbon (TOC), total alkenones and dinosterol. Shaded area indicates the deglaciation
period.
To apply marine biomarker abundances pre- served in the sediments as quantitative indicators of paleoproductivity, biomarker/cell ratios have to be constant throughout geological time. Prahl et al. (1988) reported that the alkenones:cell ratio varies with cell size of E. huxleyi with a small amplitude of 1.20_+ 0.28 pg/cell from the analyses of single- strain cultured samples, while Conte and Eglinton (1993) analyzed surface water samples collected from the eastern North Atlantic and reported that the ratios may range by about one order of magni- tude (0.6 7 pg/cell) depending on blooming state or nutrient availability. Because alkenones are bio- synthesized not only by E. huxleyi but also by other prymnesiophytes, alkenones:cell ratios may be vari- able among species. Therefore, strictly speaking, paleoproductivity of prymnesiophyte algae cannot be quantitatively estimated even from alkenone abundances in the sediments.
At this moment the best way to minimize the misinterpretation caused by secular variations of biomarker:cell ratios in estimating the biomarker- based paleobiological productivity is to analyze as large numbers of marine-derived compounds as possible and to find a common downcore trend in their profiles. This is because each compound gener- ally has different metabolic roles in the cells of source organisms and the factor(s) controlling bio- marker:cell ratio is different from one biomarker to another. For example, alkenones are thought to be
biosynthesized as one of the components of cellular membrane (Marlowe el al., 1984b). Pristane and lower molecular weight n-alkanes are synthesized as materials of buoyancy regulation, thermal insula- tion and energy storage in marine organisms (Nevenzel, 1970; Lee et al., 1971), although pristane is also derived from the degradation of phytol side- chain of chlorophyll through the oxidation pathway (Didyk et al., 1978). Steroids such as dinosterol probably act as reinforcers in the cellular membrane (Rohmer et al., 1979).
We analyzed the downcore sediments (KH92-1- 5cBX) for estimating MARs of total organic carbon (TOC) and four kinds of marine biomarkers: total LMW n-alkanes; pristane; total alkenones; and dinosterol (Fig. 6). All exhibit synchronous decrease during the deglaciation period. The lowest flux of alkenones was observed at the latest stage of the deglaciation period, which is about 1/40 of the highest wdue at the LGM. However, it is unlikely that the biological production decreased to such a low level; this variation may be attributed mainly to the changes in alkenones:cell ratio of source organ- isms. Nevertheless, a common trend exists among the MARs of TOC and four kinds of marine bio- markers, with relatively low values during the deglaciation period. Exponential diagenetic models are constructed to subtract degradation trend from the downcore profiles of LMW n-alkanes and pris- tane (Fig. 7). Assumptions are that: (1) biological
Climate variabilities
production and biomarker:cell ratio during the Holocene is almost constant: and (2) that decompo- sition rates of these compounds has not varied during the last 20 000 years. The results suggest low biological productivity (half that of the Holocene) during the deglaciation period and high (1.5-2 times that of the Holocene) during the LGM (Fig. 7).
Prahl (1992) found that during the LGM (15- 26 kyr) concentrations of alkenones and dinosterol are high in the sediments from the central tropical Pacific and interpreted this as the biological pro- ductivity being significantly high in this region during that time. Sikes and Keigwin (1994) reported total alkenone concentrations from the tropical Atlantic sediments and pointed out the higher bio- logical productivity during the LGM. Higher MARs of alkenones in our downcore sediments also indicate higher primary productivity during the LGM in the western tropical Pacific.
Sea surface temperature
Homologs of long-chain alkenones preserved in the deep-sea sediments have been one of the most
in the late Quaternary 179
intensively studied biomolecules of the last decade, with the knowledge that their unsaturation degree provides a tool for reconstruction of sea surface temperature (SST) in the past ocean (Brassell et al., 1986; Prahl and Wakeham, 1987). Long-chain (C37- C39 ) di-, tri- and tetra-unsaturated alkenones are synthesized only by prymnesiophyte algae (Volkman et al., 1980; Marlowe et al., 1984a, 1990). Unsaturation degree of C37 alkenones depends on the temperature of seawater where they are bio- synthesized (Marlowe et al., 1984a; Brassell et al., 1986). This is believed to be due to the biological regulation of the melting points of cellular mem- branes in order to maintain fluidity (Marlowe et al., 1984b). Unsaturation degree of C37 alkenones (UK37) has been defined by the following equation (Marlowe et al., 1984a; Brassell et al., 1986):
UK37 = ( [ C 3 7 : 2 ] - [C37:4])/[C37:2] ']- [ C 3 7 : 3 ] + [ C 3 7 : 4 ] )
where [Cx:y] denotes the concentration of the alke- none with x carbon numbers and y double bonds. Although temperature calibration of UK37 appears to be slightly dependent on the species which bio- synthesize alkenones (Brassell, 1993; Volkman et
0
5.
10.
15.
20.
25 0
o
(a)
. . . . ; . . . . 1~) . . . . 15 . . . . 2 0
C17-C20 n - ~ $ ( n g / k y r / e m 2 )
0 1 2 3
R e l a t i v e m a s s a c c u m u l a t i o n ra te of Ct7-C2o n - a l k a n e s
0
5.
"~ 10-
g~ < 1 5 -
20-
25
°/o (b)
. . . . I . . . . I . . . . I . . . . 1 2 3 4
Pristaae (ng/kyr/em2)
<
0 0.5 1 1.5 2 Relative mass accumulation rate
of prislane
Fig. 7. Reconstruction of marine biological productivity during the last 20 kyr. (a) and (b) Thick lines indicate exponential diagenetic models to subtract degradation effect from the downcore profile of C17- C2o n-alkanes and pristane, respectively, in KH92-1-5cBX (O). These models assume that the fluxes of these biomarkers did not vary significantly during the Holocene (0-7 kyr BP) and that the decompo- sition rates of these biomarkers are constant during the last 20 kyr. (c) and (d) The relative fluxes of total C17-C20 n-alkanes and pristane, respectively, during the last 20 kyr. They are shown as the ratios
of analytical values to the models. Shaded area indicates the deglaciation period.
180 N. Ohkouchi et al.
al. , 1995), recently Brassell (1993) compiled all the data reported since 1984 (e.g. Prahl and Wakeham, 1987; Sikes and Volkman, 1993; Conte and Eglinton, 1993) and revised the calibration equation as follows:
UK37 = 0.037 x S S T - 0.083
We estimated paleo-SST in the western tropical Pacific by using the above calibration (Ohkouchi et
al. , 1994), which suggests that the UK37-based SST was almost constant (28-29°C) over the last 20000y (Fig. 8). The core site is located at the 800 km north of New Guinea mountain range at altitudes higher than 2500 m where air temperature during the LGM was revealed to have decrease by about 6°C by several pollen studies (Fig. 8; Hope, 1976; Webster and Streten, 1978). The apparent dis- crepancy between conservative change in SST and drastic temperature drop in mountain area may be caused by the difference in the present (6°C/1000 m) and the past (LGM) adiabatic lapse ranges. To reconcile the discrepancy, we proposed that the adiabatic lapse range was higher (7-8°C/1000 m) at the LGM than at the present time (Fig. 9).
In general, most UK37-based SST at the LGM previously reported exhibit consistent results to the CLIMAP reconstruction which applied a statistical approach to the microfossil assemblage preserved in the sediments (Imbrie and Kipp, 1971; CLIMAP Project Members, 1976, 1981), although some of them exhibit lower values than those from micro- fossil assemblage (Jasper and Gagosian, 1989; Sikes and Keigwin, 1994). One recently emerged contro-
Fig. 8. Upper: air temperature during the last 20 kyr in the mountain range (altitudes higher than 2500 m) in New Guinea shown as the difference from that of the present (modified from Webster and Streten, 1978). Lower: sea surface temperature (SST) during the last 20 kyr in the western tropical Pacific estimated from UK37. It is shown both as absolute temperature scale and the difference from
the present SST.
5000 "71 - 5000 4 L \ " ....
~. 3 0 0 0 - ~ - , n o w ~ ' , , \ . .3000 ~ ~ ~1 "l '% ~ montane fo i s t ~
2000!~ ", ~ i m i t monUme forest '%, ~ .2000
1000 "-] ~ ~ .~t" , 1000
0 4 ..................... -'~.. ..... o ] . . . . i . . . . i . . . . i , , , ~ i , ,
-10 0 10 20 30 Temperature (°C)
Fig. 9. Adiabatic lapse ranges at the present (thick line) and the LGM (broken line) in the western tropical Pacific. The altitudes of the present snow line and montane forest
limit and LGM were determined by Hope (1976).
versy of Quaternary paleoclimatology is SST esti- mation in the tropical ocean at the LGM. Microfossil assemblage, UK37 and foraminiferal 61So records suggested a glacial-interglacial SST difference in the tropical ocean of at most 2°C (CLIMAP Project Members, 1981; Broecker, 1986; Prahl et al. , 1989; Sikes and Keigwin, 1994; Ohkouchi e t al. , 1994), while a recently established paleothermometer, coral Sr:Ca ratio, as well as coral 6180 record suggested that SST was lower by 5°C in the glacial tropical ocean (Aharon and Chappell, 1986; Beck et al. , 1992; Guilderson et al. ,
1994). The discrepancy between plankton- and coral-based paleotemperatures may be caused by ecological and/or physiological effects on these organisms which have been neglected.
SUMMARY
1. The latitudinal distribution of C25 C35 n-alkanes preserved in the pelagic sediments exhibits high concentrations in high latitudes (48-25°N) in the Northern Hemisphere and low in middle to low latitudes (13°N-15°S), being similar to that of present atmospheric aerosol concentrations over the Pacific Ocean. The parallel distributions between the pelagic sediments and the recent marine atmosphere support a long-range atmos- pheric transport of continentally derived HMW n-alkanes to the deep-sea sediments.
2. The downcore distribution of total Cz5-C35 n- alkanes suggests that during the last glacial maximum the atmospheric transport of terres- trial materials over the ocean was doubled com- pared with that of the Holocene, while during the last half of the deglaciation period it decreased down to half that of Holocene levels in the western tropical Pacific.
3. The latitudinal distributions of sedimentary C17- C20 n-alkanes and pristane is high in the higher latitudes in the Northern Hemisphere and
Climate variabilities in the late Quaternary 181
a r o u n d the Equator , while remain ing low in the mid- la t i tude regions. Their d is t r ibut ions p rob- ably reflect the biological product ivi ty in the ocean.
4. Downcore dis t r ibut ions of C17-C20 n-alkanes, pristane, long-chain a lkenones and dinosterol suggest tha t in the western t ropical Pacific the pr imary product ivi ty was significantly depressed dur ing the deglaciat ion per iod and was increased slightly dur ing the L G M compared with the Holocene.
5. Sea surface tempera ture deduced f rom the unsa- tu ra t ion degree of a lkenones was a lmost cons tan t dur ing the last 20 000 y in the western tropical Pacific. Coupled with pollen results f rom New Guinea m o u n t a i n range suggests tha t the adia- batic lapse range in New Guinea was higher (7 - 8°C/1000 m) than tha t of the present day (6°C/ 1000 m).
Acknowledgements--We appreciate the encouragement and useful comments of Prof. R. Ishiwatari. We would also like to thank Drs. P. Hesse, H. Kawahata, I. Koike, R. Matsumoto, T. Oba, R. Tada, and M. Uematsu for their constructive comments to our results. Thanks are also due to Profs. S.A. Macko and H.R. Harvey for pro- viding us a chance to present this paper. This study was partially supported by the grant from Japan Society for the Promotion of Science to N.O. American Chemical Society financially supported N.O.'s stay at Penn State University.
REFERENCES
Aharon, P. and Chappell, J. (1986) Oxygen isotopes, sea level changes and the temperature history of a coral reef environment in New Guinea over the last 105 years. Paleogeogr. Paleoclimatol. Paleoecol. 56, 337-379.
Barnola, J. M., Raynaud, D., Korotkevich, Y. S. and Lorius, C. (1987) Vostok ice core provides 160 000-year record of atmospheric CO2. Nature 329, 408-414.
Beck, J. W., Edwards, R. L., Ito, E., Taylor, F. W., Recy, J., Rougerie, F., Joannot, P. and Henin, C. (1992) Sea- surface temperature from coral skeletal strontium/cal- cium ratios. Science 257, 644-647.
Berger, W. H. (1989) Global maps of ocean productivity. In Productivity o f the Ocean: Present and Past (Edited by Berger, W. H., Smetacek, V. S. and Wefer, G.), pp. 429-455. John Wiley, Chichester.
Berner, R. A. (1980) Early Diagenesis: A Theoretical Approach. Princeton University Press, Princeton.
Betzer, P. R., Carder, K. L., Duce, R. A., Merrill, J. T., Tindale, N. W., Uematsu, M., Costello, D. K., Young, R. W., Feely, R. A., Breland, J. A., Bernstein, R. E. and Greco, A. M. (1988) Long-range transport of giant min- eral aerosol particles. Nature 336, 568-571.
Boon, J. J., Rijpstra, W. I. C., de Lange, F., de Leeuw, J. W., Yoshioka, M. and Shimizu, Y. (1979) Black Sea sterol--a molecular fossil for dinoflagellate blooms. Nature 277, 125-128.
Brassell, S.C. (1993) Applications of biomarkers for deli- neating marine paleoclimatic fluctuations during the Pleistocene. In Organic Geochemistry: Principles and Applications (Edited by Engel, M. H. and Macko, S. A.), pp. 699-738. Plenum Press, New York.
Brassell, S. C., Eglinton, G., Marlowe, I. T., Pflaumann, U. and Sarnthein, M. 0986) Molecular stratigraphy: a new tool for climatic assessment. Nature 320, 129-133.
Broecker, W. S. (1986) Oxygen isotope constraints on sur- face ocean temperatures. Quat. Res. 26, 121-134.
Broecker, W. S. and Sarmiento, J. (1980) GEOSECS Pacific Ocean Expedition, Government Printing Office.
Budzikiewicz, H., Djerassi, C. and Williams, D. H. (1967) Mass Spectrometry of Organic Compounds. Holden-Day, New York.
CLIMAP Project Members (1976) The surface of the ice- age earth. Science 191, 113 l- 1137.
CLIMAP Project Members (1981) Seasonal reconstruc- tions of the earth's surface at the last glacial maximum, Geol. Soc. Am. Map Chert MC-36, 1-18.
Conte, M. H. and Eglinton, G. (1993) Alkenone and alkenoate distributions within the euphotic zone of the eastern North Atlantic: correlation with production tem- perature. Deep-Sea Res. 40, 1935-1961.
Didyk, B. M., Simoneit, B. R. T., Brassell, S. C. and Eglinton, G. (1978) Organic geochemical indicators of paleoenvironmental conditions of sedimentation. Nature 272, 216-222.
Duce, R. A. (1989) SEAREX: The sea-air exchange pro- gram. In Chemical Oceanography 10 (Edited by Riley, J. P. and Chester, R.), pp. 1-14. Academic Press, New York.
Duce, R. A., Unni, C. K., Ray, B. J., Prospero, J. M. and Merrill, J. T. (1980) Long-range atmospheric transport of soil dust from Asia to the tropical North Pacific: temporal variability. Science 209, 1522-1524.
Engel, M. H. and Macko, S. A. (ed.) (1993) Organic Geochemistry. Plenum Press, New York.
Eglinton, G. and Calvin, M. (1967) Chemical fossils. Sci. Am. 216, 32-43.
Eglinton, G. and Hamilton, R. J. (1967) Leaf epicuticular waxes. Science 156, 1322-1335.
Eglinton, G., Bradshaw, S. A., Rosell, A., Sarnthein, M., Pflaumann, U. and Tiedemann, R. (1992) Molecular record of secular sea surface temperature changes on 100-year timescales for glacial terminations I, II, and IV. Nature 356, 423-426.
Gagosian, R. B. and Peltzer, E. T. (1986) The importance of atmospheric input of terrestrial organic material to deep sea sediments. Org. Geochem. 10, 661-669.
Graccum, R. A. and Prospero, J. M. (1980) Saharan aero- sols over the tropical North Atlantic--mineralogy. Mar. Geol. 37, 295-321.
Guilderson, T. P., Fairbanks, R. G. and Rubenstone, J. L. (1994) Tropical temperature variations since 20000 years ago: modulating interhemispheric climate change. Science 263, 663-665.
Harwood, J. L. and Russell, N. J. (1984) Lipids in Plants and Microbes. Allen and Unwin, London.
ten Haven, H. L., Baas, M., Kroot, M., de Leeuw, J. W., Schenck, P. A. and Ebbing, J. (1987) Late Quaternary Mediterranean sapropels. III: assessment of source of input and palaeotemperature as derived from biological markers. Geochim. Cosmochim. Acta 51, 803-810.
Hooghiemstra, H., Bechler, A. and Beug, H.-J. (1987) Isopollen maps for 18000 years B.P. of the Atlantic off- shore of northwest Africa: evidence for paleowind circu- lation. Paleoceanogr. 2, 561-582.
Hope, G. S. (1976) The vegetation history of Mt Wilhelm, Papua, New Guinea. J. Ecol. 64, 627-664.
Ikehara, M. (1994) Paleoceanographic changes in the wes- tern equatorial Pacific Ocean during the last 200 kyrs. MS Thesis, Kanazawa University.
Imbrie, J. and Kipp, N. G. (1971) A new micropaleontolo- gical method for paleoclimatology: Application to a Late Pleistocene Caribbean core. In The Late Cenozoic
182 N. Ohkouchi et al.
Glacial Ages (Edited by Turekian, K.K.), pp, l-Sl. Yale University Press, New Haven.
Ishiwatari, R., Hirakawa, Y., Uzaki, M., Yamada, K. and Yada, T. (1994) Organic geochemistry of the Japan Sea sediments 1: bulk organic matter and hydrocarbon analyses of core KH79-3, C-3 from the Oki Ridge for paleoenvironment assessments. J. Oceanogr. 50, 179 195.
Janecek, T. R. and Rea, D. K. (1985) Quaternary fluctu- ations in northern hemisphere tradewinds and westerlies. Quat. Res. 24, 150-163.
Jasper, J. P. and Gagosian, R. B. (1989) Alkenone mol- ecular stratigraphy in an oceanic environment affected by glacial freshwater events. Paleoceanogr. 4, 603 614.
Johns, R. B. (ed.) (1986) Biological Markers in the Sedimentary Record. Elsevier, Amsterdam.
Kawamura, K. (1995) Land-derived lipid class compounds in the deep-sea sediments and marine aerosols from North Pacific. In Biogeochemical Proeesses o/" Ocean Flux in the Western Pacific (Edited by Sakai, H. and Nozaki, Y.), Terra Sci. Publ., in press.
Kawamura, K. and lshiwatari, R. (1985) Distribution of lipid-class compounds in bottom sediments of fresh- water lakes with different trophic status. Chem. Geol. 51, 123 133.
Kennedy, J. A. and Brassell, S. C. (1992) Molecular records of twentieth-century El Nifio events in lami- nated sediments from the Santa Barbara basin. Nature 357, 62-64.
Lee, R. F., Hirota, J. and Barnett, A. M. (1971) Distribution and importance of wax esters in marine copepods and other zooplankton. Deep-Sea Res. 18, 1147 1165.
Marlowe, I. T., Brassell, S. C., Eglinton, G. and Green, J. C. (1984a) Long chain unsaturated ketones and esters in living algae and marine sediments. Org. Geochem. 6, 135 141.
Marlowe, I. T., Green, J. C., Neal, A. C., Brassell, S. C., Eglinton, G. and Course, P. A. (1984b) Long-chain (n- C37-C39 ) alkenones in the Prymnesiophyceae-distri- bution of alkenones and other lipids and their taxo- nomic significance. Br. Phycol. J. 19, 203 216.
Marlowe, l. T., Brassell, S. C., Eglinton, G. and Green, J. C. (1990) Long-chain alkenones and alkyl alkenoates in living algae and marine sediments. Chem. Geol. 88, 349 375.
Martin, J. H. (1990) Glacial-interglacial CO2 change: the iron hypothesis. Paleoceanogr. 5, 1 13.
Merrill, J. T. (1989) Atmospheric long range transport to the Pacific Ocean. In Chemical Oceanography 10 (Edited by Riley, J. P. and Chester, R.), pp. 15 50. Academic Press, New York.
Neftel, A., Oeschger, H., Staffelbach, T. and Stauffer, B. (1988) CO2 record in the Byrd ice core 50000 5000 years BP. Nature 331, 609 611.
Nevenzel, J. C. (1970) Occurrence, function, and biosyn- thesis of wax esters in marine organisms. Lipids 5, 308 319.
Ohkouchi, N. (1995) Lipids as biogeochemical tracers in the late Quaternary, PhD Thesis, University of Tokyo.
Ohkouchi, N., Kawamura, K. and Taira. A. (1994) Small changes in the sea surface temperature during the last 20000 years: molecular evidence from the western tropi- cal Pacific. Geophys. Res. Lett. 21, 2207-2210.
Ohkouchi, N., Kawamura. K. and Taira, A. (1997a) Organic geochemistry of carbonate-rich deep sea sedi- ments from the western tropical Pacific: Fluctuations of marine and terrestrial input during the last 20000 years. Paleoceanogr., 12, 623-630.
Ohkouchi, N., Kawamura, K., Kawahata, H. and Taira, A. (1997b) The distribution of lipid class compounds in the deep sea surface sediments: A latitudinal transect at
175E. Submitted to Geochim. Cosmochim. Acta. 61, 1911 1918.
Okada, H. and Honjo, S. (1973) The distribution of ocea- nic coccolithophorids in the Pacific. Deep-Sea Res. 20, 355 374.
Philp, R. P. (1985) Fossil Fuel Biomarkers: Applications and Spectra. Elsevier, Amsterdam.
Poynter, J. G., Farrimond, P., Robinson, N. and Eglinton, G. (1989) Aeolian-derived higher plant lipids in the marine sedimentary record: links with paleocli- mate. In Modern and Past Patterns gf Global Atmospheric TramT)ort (Edited by Leinen M. and Sarnthein, M.), pp. 435 462. Kluwer, Amsterdam.
Prahl, F. G. (1992) Prospective use of molecular paleon- tology to test for iron limitation on marine primary pro- ductivity~ Mar. Chem. 39, 167-185.
Prahl, F. G. and Wakeham, S. G. (1987) Calibration of unsaturation patterns in long-chain ketone compositions for paleotemperature assessment. Nature 330, 367--369.
Prahl, F. G., Muehlhausen, L. A. and Zahnle, D. L. (1988) Further evaluation of long-chain alkenones as indicators of paleoceanographic conditions. Geochim. Cosmochhn. Acta 52, 2303 2310.
Prah[, F. G., Muehlhausen, L. A. and Lyle, M. (1989) An organic geochemical assessment of oceanographic con- ditions at MANOP Site C over the past 26000 years. Paleoceanogr. 4, 495-510.
Prospero, J.M. (1981) Eolian transport to the world ocean. In The Sea (Edited by Emiliani, C.), Vol. 7, pp. 80l 874. John Wiley, Chichester.
Rea. D. K. (1994) The paleoclimatic record provided by eolian deposition in the deep sea: the geologic history. Rev. Geophys. 32, 159 195.
Rohmer, M., Bouvier, P. and Ourisson, G. (1979) Molecular evolution of biomembranes: structural equivalents and phylogenetic precursors of sterols. Proc. Nat. Acad. Sci. 76, 847 851.
Rostek, F., Ruhland, G.. Bassinot, F. C.. Mtiller, P. J., gabeyrie, k. D., Lancelot, Y. and Bard, E. (1993) Reconstructing sea surface temperature and salinity using 6180 and alkenone records. Nature 364, 319 321.
Sarnthein, M., Tetzlaff, G., Koopmann, B., Wolter, K. and Pflaumann, U. (1981) Glacial and interglacial wind regimes over the eastern subtropical Atlantic and north- west Africa. Nature 293, 193 196.
Sikes, E. L. and Volkman, J. K. (1993) Calibration of alkenone unsaturation ratios UK37"s for paleotempera- ture estimation in cold polar waters. Geochim. Cosmochim. Acta 57, 1883 1889.
Sikes, E. L. and Keigwin, L. D. (1994) Equatorial Atlantic sea surface temperature for the last 30 kyr: a compari- son of Uk37', 0180 and foraminiferal assemblage tem- perature estimates. Paleo~eanogr. 9, 31 45.
Simoneit, B. R. T. (1977) The Black Sea, a sink for terri- genous lipids. Deep-Sea Res. 24, 813 830.
Simoneit, B. R. T. (1978)The organic chemistry of marine sediments. In Chemical Oceanography (Edited by Riley, J.P. and Chester, R.), Vol. 7, Ch. 39, pp. 233 311. Academic Press, New York.
Tissot, B. P. and Welte, D. H. (1984) Petroleum Formation and Occurrence, 2rid edn. Springer-Verlag, Heidelberg.
Tulloch, A. P. (1976) Chemistry of waxes of higher plants. In ChemistJ3, and Biochemisto' o f Natural Waxes (Edited by Kolattukudy, P.E.), pp. 236 287. Elsevier, Amsterdam.
Uematsu, M., Duce, R. A. and Prospero, J. M. (1985) Deposition of atmospheric mineral particles in the North Pacific Ocean. J. Atmos. Chem. 3, 123-138.
Volkman, J. K., Eglinton, G., Corner, E. D. S. and Sargent, J. R. (1980) Novel unsaturated straight chain C3v-C39 methyl and ethyl ketones in marine sediments and a coccolithophore Emiliania huxleyi. In Advances in
Climate variabilities in the late Quaternary 183
Organic Geochemistry 1979 (Edited by Douglas, A. G. and Maxwell, J. R.), pp. 219-227. Pergamon, Oxford.
Volkman, J. K., Barrett, S. M., Blackburn, S. I. and Sikes, E. L. (1995) Alkenones in Gephyrocapsa ocea- nica: implications for studies of plaeoclimate. Geochim. Cosmochirn. Acta 59, 513-520.
Webster, P. J. and Streten, N. A. (1978) Late Quaternary ice age climates of tropical Australasia: interpretations and reconstructions. Quat. Res. 10, 279-309.
Zahn, R. and Pedersen, T.F. (eds) (1994) Carbon cycling in the Global Ocean: Constraints on the Ocean's Role in Global Change, Springer-Verlag, Heidelberg.
